Chapter 6 Tsunami Warning Systems Past, Present, and Future

Tsunami warning systems of the past and present are based on seismographic and tide-gage information. Recently it has been realized that tsunami detection and warning might be based on the effects created in the ionosphere by a propagat- ing long wave such as tsunami. In the first three sections the instrumentation, warning systems, and sociological problems will be discussed and in the final three sections the concept of a future tsunami warning service based on atmospheric detection of tsunamis will be included.

6.1 Seismic and Tsunami Instrumentation

SEISMICINSTRUMENTATION The basic principle of a seismograph and then some recent developments will be considered. In this discussion, the books by Richter (1958) and Hodgson (1964) are followed closely. Although seismic waves lose much of their energy during propagation, they may still be detected by a sensitive instrument called a seismometer. When a recording system is added to the seismometer it becomes a seismograph. Thus, a seismograph is an instrument that records the nature of the earth’s movement caused by an earthquake or any other disturbance. The most general movement consists of translation, rotation, and strain. Our interest is in the translation movement due to an earthquake at a distance, and rotational effects that may be important only close to the epicenter (Bullen 1963). The earliest seismograph appears to have been invented in China in the year 132. This instrument consisted of a ring in which several dragon mouths, each holding a ball, were arranged. When an earthquake occurred, depending on the direction of movement, a particular ball would drop into the mouth of a dragon. By mid-19th century, Italy had seismographs that recorded the time of the earthquake by making a mark on a rotating chronograph drum. Present seismographs, based on the principle of the pendulum, were developed in Japan around 1880; then in Italy, and later in other European countries. The principle of a pendulum-type seismograph is: in an ideal situation a heavy mass, because of its inertia, would not move at all, while the ground would move following the earthquake. As this cannot be realized in practice, the mass is restrained by a spring or some other restoring force, to return the mass to its original position relative to the ground. The spring should not be strong enough to alter the motion significantly.

28 1 FIG. 6.1. Schematic diagram of a horizontal pendulum, the simplest form of seismometer. (Hodgson 1964)

The horizontal pendulum is shown in Fig. 6.1. A large mass, M, is held in position by a wire, AM, and a strut, MB, pivoted at B; A and B are points near the top and bottom, respectively, of a vertical support, AB (Hodgson 1964). By em- bedding this vertical support in a concrete pier in the ground a seismometer may be obtained. In the ideal case of zero friction at B, when an earthquake occurs the pier and the upright support will move in response to the incoming seismic waves; however, the inertia of M will tend to keep it from moving. The differential motion between M and the pier is the earthquake signal at this seismometer. In practice, improvements had to be made to this simple system. Once the pendulum starts to swing because of an earthquake, it will continue to do so, and its motion as time goes on will be unrelated to the actual ground movement. Hence, the pendulum must be stopped after it has recorded the initial ground motion. There are several ways of doing this; a paddle attached to the strut, BM, moves in a pool of oil, which will quickly damp the motion, or alternately a copper vane on the strut moves between the pole pieces of a magnet. The motion of the strut in the vane sets up eddy currents that oppose the motion and thereby bring the pendulum to rest. The motion can be recorded in several different ways: a pen attached to the strut would draw a straight line on a moving paper roll, but when an earthquake occurs this line would zigzag. A system of levers could be used to magnify the motion of the strut from io2 to io5 times so that distant earthquakes could be recorded. A mirror and light arrangement for magnification with recording on photographic paper is an alternative. At present it is common to record electromag-

282 netically by attaching a coil to the strut and making this coil move in a magnetic field, and thereby generate a current, operate a galvanometer on this current, and record on a moving paper roll either by pen and ink or photographically. To conserve paper, it is wound round a rotating drum and the paper moves sideways as it rotates, so when it is removed from the drum, the record appears in the form of a number of parallel lines. Knowledge of arrival time of different waves at the station is necessary to determine the time of the earthquake. This can be obtained by providing a time scale, as marks at every minute or so, on the record. Because these times must be accurate, crystal clocks are used at present. The site for a seismograph has to be selected carefully because extraneous factors such as microseisms might mask the true record. It is advisable to locate the seismograph in hardcore rock rather than in soft ground. To describe the translational movement of the ground due to the earthquake three separate components have to be measured; the east-west component, the north-south component, and the vertical component. Thus, at a given seismographic station a complete set of seismographs consists of two horizontal seismographs and a vertical seismograph. Seismic waves have periods ranging from a fraction of a second to several minutes, and no single instrument can give a uniform response over such a range. Thus, different instruments are needed for different ranges: short-period seismographs are used for the period range of 0.2-2 s, and long-period seismographs for the 15-100 s, intermediate seismographs cover the range of 2-15 s, and ultralong- period instruments for periods over a minute. Near earthquakes, with predominant periods of a fraction of a second, could be recorded by short-period seismographs whereas distant earthquakes, usually with surface waves with periods of more than 2 s, could be recorded through long-period seismographs. Bullen (1963) mentioned that if the station has only one instrument, usually a seismograph with good response over the range 6-8 s is used. The important observatories contain three seismographs to cover three ranges. If the observatory is located in a seismically active region, then it has, in addition to these, a low-sensitive, strong-motion seismograph that cannot be knocked out by a strong earthquake locally. / The following three parameters specify a simple seismometer (Richter 1958): the free period, the damping constant, and the static magnification. Another parameter called dynamic magnification is the amount by which the magnitude of the ground motion is magnified. A seismometer of free period much shorter than the periods of the waves it records is called an accelerometer, whereas in the reverse situation, it is called a displacement meter. In the presence of damping, the free period, 7,is modified to r/j where j is a constant that satisfies the relation h2 + j2 = 1, h being an instrumental constant. To determine 7, the damping part of the instrument is removed and the free oscillations are timed, or the period of a damped oscillation is timed and corrected. Although the instrumental constant, h, could be used to specify the damping, in practice another constant, E, related to h through E = e-nh/i, is used. Thus E is the

283 ratio of two successive swings (in opposite directions) of the motion given by h2 + j2 = 1. The static magnification, V, of the instrument can be understood if T is the period and A is the amplitude of the simple harmonic motion in the ground at the seismograph station. If the pendulum performs an oscillation with period, T, and amplitude, B, then the seismogram gives a period, T, and amplitude, VB. Until the beginning of the 1920s, the important consideration in seismograph design was to lengthen the period of the pendulum so that the dynamical magnification was nearly uniform. According to the formula T = 2n which gives the period, T, of a simple pendulum of length, L, under the action of gravity, g, a pendulum with a period of 10 s needs a cumbersome length of 25 m. In the Milne-Shaw seismograph, a period of 12 s was achieved by designing the instrument so that only a small part of g is used to control a horizontal bracket pendulum. The restoring force for the inverted pendulum is a heavy mass balanced on a knife edge coupled with springs. Recording is done photographically on smoked paper. A seismograph weighing 22,353 kgm was built in Zurich (Richter 1958). Actually, the tiny torsion seismometers developed by Anderson and Wood in 1922 are much more effective than these cumbersome seismographs. Bilham and King (1970) discussed the recent advances in strainmeters and compared their advantages and disadvantages with other geophysical instruments. They considered the frequency response and amplitude response of strainmeters. Laser interferometers were also included in their discussion. For recent advances in seismic instrumentation, see Plesinger (1970), Vogel (1970), Willmore (1970), Berckhemer (1970), Husebye (1970), Bonjer and Mueller (1970), Husebye and Bungum (1 970), Weichert (1970), and Smith (1975). A submarine seismograph developed by Koresawa ( 1963) will be described briefly. As shown in Fig. 6.2, this system consists of an electrodynamic transducer of vertical component, a transistorized amplifier, an oscillograph, a clock, a galvanometer, and a photographic recording apparatus driven by a small motor. This whole assembly is enclosed in an iron case to which heavy ballasts (to put the

Galvanomstar

lime Switch

Benery111 I I6V

1FIG.6.2. Submarine seismograph. (Koresawa 1963)

284 seismograph at the ocean bottom) are attached, as well as a float and an arrangement for refloating the seismograph after the ballasts are detached. When the observation is made and the instrument rises to the sea surface, an alarm attached to the system serves as a warning. Nagumo et al. (1965) described the details of such an ocean-bottom seismograph. Kishinouye (1966) described a modification to the earlier ocean-bottom seismograph. In the latter version, not only the vertical motion at the sea bottom, but also the other components (east-west and north-south) of the motion could be recorded. Ewing and Ewing (196 1) described a telemetering ocean-bottom seismograph that will rest either on the ocean bottom or planted in the sediments. Arnett and Newhouse (1965) described an ocean-bottom seismograph developed as a part of the project “vela uniform” at the Air Force Cambridge Research Laboratories by Texas Instruments. This was used to study the acoustic wave energy generated either by earthquakes or by explosions at the interface of the solid earth with the water. Another important paper dealing with ocean-bottom seismographs is that of Sutton et al. (1965), who described seismic and geophysical measurements at the ocean bottom through seismic observatories. Bradner et al. (1965, p. 1906) discussed the reliability of seismic measurements made at the ocean bottom and described a technique of improving the reliability. “Motion of the seismometer cases and seismometers due to ocean currents, thermal convection, or mechanical motions of parts in the seismometer package could be responsible for details in the apparent seismic spectra. Thermal convection can be caused, for example, by heat dissipation in the instrument electronics. The validity of the spectra can be most easily checked by making detailed digital analyses of simultaneous spectra from two proximate three-component ocean-bottom instruments.” TSUNAMIGAUGES The Hawaiian Institute of Geophysics has developed sophisticated gauges to record tsunamis in the deep ocean. Some important papers are Vitousek (1963b), and Vitousek and Miller (1970). It was claimed that tsunamis in the open ocean at depths to 6 km could be measured. Filloux (1970a, b) described Bourdon tube deep-sea gauges. Homma et al. (1966) described the response characteristics of underwater gauges. Benioff and Gutenberg (1939) described a portable tsunami recorder; Snodgrass (1958) discussed a shore-based recorder of low-frequency ocean waves; Van Dorn (1960a) described a long-period wave recorder for the range of 10-I05s; and Larsen (197 1) discussed the electromagnetic field of long and intermediate water waves.

MICROBARAGRAPHSAND AEROLOGICALDATA FOR TSUNAMIPREDICTION The aerological aspects of the tsunami and related problems will be discussed in subsequent sections. Some microbarographs that could be used to record pressure fluctuations in the atmosphere due to any perturbation will be briefly described. Van Dorn ( 1960b) described a low-frequency microbarograph that was used by the Scripps Institution of Oceanography, La Jolla, Calif. A hydraulic

285 resistance-capacitance filter network attenuates the high-frequency gusts and atmospheric tide. The resolution of the instrument was f 0.05 mb and the instrument could attain sensitivities up to 0.1 mb. Donn (1958) described a microbarovariograph with a maximum resolution of 0.002 mb. Hines (1972, p. 78) proposed that atmospheric internal gravity waves might be used for tsunami prediction. “Tsunamis have their origins in major earthquakes beneath the ocean floor or at its verge. A tsunami rises to rampaging amplitudes only in shallow waters, typically at the coasts themselves, where its energy must be carried by relatively small masses of water. In the open ocean, a tsunami is very tame indeed, producing, perhaps, a few meters’ rise or fall of water, and that spread over a horizontal peak-to-trough distance of many km. Tsunamis are therefore virtually undetectable in the open ocean, and the prediction of their appearance on shore is most difficult. But a tsunami must displace the atmosphere as it propagates and the displaced atmosphere must respond by generating a gravity wave. The parameters are such that these waves will be of the internal type, and so will grow exponentially with height. A rise of a few meters at the surface of the water might well amplify to a few km at ionospheric heights, and that sort of amplitude could hardly escape detection if it were sought. We arrive then, at this speculative question: if we wish to keep track of the progress of a tsunami, and so predict with some assurance the onslaught of its destructive force, might we not serve our interests best by keepingwatch on the ionosphere?’ One interesting phenomenon to briefly consider is the biological precursors to tsunamis. Musya (1934), Terada (1934), and Adams (1974) discussed the luminous phenomena associated with the Sanriku earthquake tsunami of 1933. Adams ( 1974, p. 65) states: “There haye been numerous reports on lights accompanying earthquakes. Since these reported lights were usually observed concomitant with the earthquakes, they are of no predictive value. “In the case of tsunamis, the light seems to often occur from the direction of the sea. One possible explanation is that the light is generated by luminiscent plankton excited by the tsunami. If this explanation is correct, then the light provides a natural signal of the impending tsunami (Terada, 1934) . . . other observers of light associated with a tsunami attributed it to luminiscence of the animals on the sea bottom, exposed during ebbing of the tsunami. This would be of value to a warning system only if the tsunami began as a withdrawal.” Adams mentions that the 1933 tsunami-earthquake pair exhibited anomalous behavior of fishes several hours before the event and gave two specific examples (surface fish feeding on ocean-bottom diatoms and migration of deepwater fish).

6.2 Tsunami Protection Measures It is obvious there is no foolproof method to protect against tsunami damage to property; nevertheless, breakwaters might offer some protection. There is an enormous amount of literature on breakwaters for protection against short-period waves, but it will not be included in this Bulletin. NUMERICALEXPERIMENTS It6 (1970, 1971) gave some interesting results on tsunami waves in Ofunato Harbor with and without a breakwater (Fig. 6.3). The tsunamis of 1896, 1933,

286 m 2.5 - -With Breakwatsr 2.0 -

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and 1960 caused heavy damage in Ofunato Bay. A breakwater was constructed in 1967 and a year later, another tsunami occurred. Thus, this case provides an opportunity to examine tsunami waves with and without breakwater. Ito used a numerical model to compute the amplitude of the tsunami inside the harbor with and without a breakwater. Figure 6.3 shows that the breakwater indeed damps the tsunami inside the harbor. Murty and Boilard (1970) calculated the effect of a hypothetical breakwater on the tsunami amplitude in Alberni Inlet on the west coast of Canada and their results are in general agreement with those of Ito. LABORATORYEXPERIMENTS Kamel ( 1970) described hydraulic model experiments specifically to understand design criteria for breakwaters to protect Hilo from incoming tsunamis. The experiments were performed in a steel flume 148.1 m long, 1.8 m wide, and 1.2 m deep. The ratio of the model to prototype was 1:50. Some preliminary experiments were made in a three-dimensional model to simulate the tsunami bore, and the data obtained on the height and shape of the bore as well as the height and duration of the overtopping above the barrier were used in the flume experiments. Based on these experiments, Kamel concluded that efforts to develop an overtopping rubble-mound tsunami barrier with a steep harborside slope were not successful. Wiegel (1960) studied theoretically and through laboratory experiments the wave transmission past a thin rigid barrier, extended above the water surface and to some distance below the water surface. The aim of Wiegel's study was to determine the depth to which such a barrier (realized in practice by mounting a siding on a pile structure) had to be immersed for effectively stopping the waves.

287 Wiegel concluded that a rigid barrier extending only a short distance below the surface is not useful as a protection measure.

FORCEOF TSUNAMISURGE ON A STRUCTURE Fukui et al. (1963) showed that the maximum pressure, p,,,, by a tsunami on a vertical (or nearly vertical) wall is

where K is a constant empirically determined to be 0.5, Y is the specific weight of water, g is gravity, D is the height of the bore, and u is the velocity of the tsunami advancing as a bore over dry bed and is given by :

u = 1.83 (6.2) The pressure is maximum at the bottom of the wall and decreases linearly to zero at a height of 1.5 D. Matlock et al. (1962) studied the structural damage at Hilo due to the 1960 Chilean earthquake tsunami. The pressure on the structures was estimated as the sum of the hydrostatic pressure and the dynamic pressure, pD,given by :

where p is the density of water and the drag coefficient, C,, was taken as 1.2. Cross (1967) studied, both theoretically and through laboratory experiments, tsunami surge force due to the propagation of tsunami as a bore over drybed. Cross assumed that the shape of the front portion of the tsunami bore is determined mainly by friction and hydrostatic pressure gradient induced by the surface slopes. Another assumption is that the velocity and acceleration at the front are mainly determined by factors external to the region of the tip of the surge and these were assumed to be known. The present problem is in several respects very similar to the turbidity current problem discussed in Chapter 2. Cross used momentum considerations to arrive at the following relation for the force, F, on the wall:

F = 'YD2 2 + C,pu*D (6.4) where C, is a force coefficient defined by:

FD c, = pu2 D (6.5) where F, is the dynamic force on the wall. Equations (6.4), (6.5) can be used U to calculate the force exerted by the tip of the tsunami surge on a vertical wall. Cross's laboratory experiments agreed fairly well with his theoretical results. The discrepancies were attributed to the neglect of convective accelerations in the tip region.

288 6.3 Tsunami Warning Systems of the Past and Present It is difficult to allocate a date when the idea of tsunami warning systems was conceived. Following the Aleutian earthquake of Feb. 3, 1923, Finch (1924) suggested the possibility of tsunami prediction for warning purposes. However, he stated that his idea was not new and as early as 1904 speculation was made on the possibility of warning against tsunamis. Before the development of seismic instruments, the occurrence of a major earthquake was taken as the first indication of a tsunami. For a continuously evolving problem such as tsunami warning service, any descriptions of the existing system are likely to become out-dated. Nevertheless, in the following pages, how the tsunami warning service came into existence and how it functions at present will be described.

THEUS. COASTAND GEODETICSURVEY’S TSUNAMI WARNING SYSTEM Following the disastrous Aleutian earthquake tsunami of Apr. 1, 1946, which killed 173 people and caused property damage of about $25 million in Hawaii, a group of scientists in the US. Coast and Geodetic Survey conceived the idea of a tsunami warning service for the Pacific Ocean by seismographs to detect the earthquakes, and tide gages to observe the tsunami. However, in 1946, seismographic recording was done photographically and photographs were not developed until the following day. Other difficulties were: neither the tide gages were suitable to distinguish between tsunamis and any other type of wave, nor were there any communication circuits for rapid transmission of messages. By 1948, a functional tsunami warning service was achieved. Seismographs with continuous and visible records were developed and installed at Hawaii and other stations in the Pacific. These were equipped with automatic alarm signals. A seismic seawave detector (that could filter out tides and wind waves but leave the tsunami) had been developed and installed at several locations. Better tsunami travel-time charts were prepared with accuracy of k 1.5 min for every hour of travel. Communication channels were established through the Federal Aviation Agency (at that time it was referred to as the Civil Aeronautics Administration) and through U.S. military services. Later National Aeronautics and Space Administration (NASA) and the U.S. Weather Bureau joined the communication network. The Seismic Sea Wave Warning System (SSWWS) started operating in 1948 at the Honolulu Magnetic and Seismological Observatory, where a 24-h watch was maintained on large earthquakes and tsunamis. Although originally it was intended to provide tsunami warning for Hawaii, later its role was expanded to include several other nations bordering the Pacific. Figure 6.4 shows the seismograph stations and tide-gage stations on which SSWWS was based. When an earthquake of magnitude large enough to generate a tsunami occurs, the automatic seismic alarm of each seismograph station responds at once and the seismograph stations report arrival times of P and S waves to Honolulu. Based on this data, Honolulu determines the epicenter and if the focus falls in or near the ocean, it issues an advisory bulletin. This bulletin is given

289 FIG. 6.4. Tsunami warning system showing seismic seawave travel times to Honolulu; tidal (o), and seismographic (A) stations. (International Tsunami Information Center Honolulu) to SSWWS participants and contains information on the earthquake epicenter and the possibility of tsunami generation. Also, estimated time of arrival (ETA) to each participating station is given, based on the tsunami travel-time curves. The next step in the SSWWS involves interrogation of the tide stations, especially those closer to the epicenter, to detect any unusual activity, in which case Honolulu issues a tsunami warning. However, local warning and evacuation of people is up to the designated agencies of the participating countries. The chronological events associated with the SSWWS activities during the 1964 Alaskan earthquake tsunami are taken from Anon. (1965, p. 13-24). 0336:13 (Greenwich Mean Time, March 28, 1964) Earthquake strikes the northern shore of Prince William Sound, Alaska. This disturbance at 8.5 on the Pasadena scale (Richter scale) of earthquake magnitude is the most severe ever reported for North America. (In Alaska, where it is Good Friday afternoon, the earthquake is called the Good Friday Earthquake by the press). The warning alarm sounds at Honolulu Observatory and at four other participating seismograph stations. The SSWWS begins its race with time. 0344. Seismic seawave warning alarm sounds at Honolulu Observatory. Re- quests issued for immediate readings from seismograph stations at College, Alaska; Sitka, Pasadena, Berkeley, Tucson, Tokyo, and Guam.

290 0452 Honolulu Observatory completes preliminary determination of location of earthquake epicenter: 61N, 147.5W. 0502 Honolulu Observatory issues first advisory via FAA and Defense Communications network: THIS IS A TIDAL WAVE ADVISORY. A SEVERE EARTHQUAKE HAS OCCURRED AT LAT. 61N, LONG. 147.5W, VICINITY OF SEWARD, ALASKA, AT 03362,28 MAR. IT IS NOT KNOWN, REPEAT NOT KNOWN, AT THIS TIME THAT A SEA WAVE HAS BEEN GENERATED, YOU WILL BE KEPT INFORMED AS FURTHER INFOR- MATION IS AVAILABLE. IF A WAVE HAS BEEN GENERATED, ITS ETA FOR THE HAWAIIAN ISLANDS (HONOLULU) IS 09002,28 MARCH.. . 0530 Communications inoperative on Alaskan mainland. Honolulu Observa- tory issues second bulletin. This is an information bulletin, stating that it is not yet known whether a seismic sea wave has been generated, but providing all participants in the SSWWS with estimated times of arrival. The observatory also requests inspection of tide records by observers at Unalaska, Kodiak, Adak, Sitka, Alaska; and Crescent City, Calif. 0555 Kodiak replies: EXPERIENCE SEISMIC SEA-WAVE AT 04352. WATER LEVEL 10-12 FT ABOVE MEAN SEA LEVEL. WILL ADVISE. 0630 Another message from Kodiak tide observer confirms existence of seismic sea wave. Honolulu Observatory issues third bulletin: THIS IS A TIDAL WAVEISEISMIC SEA-WAVE WARNING. A SEVERE EARTH- QUAKE HAS OCCURRED AT LAT. 61N, LONG. 147.5W, VICINITY OF SEWARD, ALASKA, AT 03362, 28 MAR. A SEA WAVE HAS BEEN GENERATED WHICH IS SPREADING OVER THE PACIFIC OCEAN. THE ETA OF THE FIRST WAVE AT OAHU IS 09002,28 MARCH. THE INTENSITY CANNOT, REPEAT, CANNOT BE PREDICTED. HOWEVER, THIS WAVE COULD CAUSE GREAT DAMAGE IN THE HAWAIIAN ISLANDS AND ELSEWHERE IN THE PACIFIC AREA. THE DANGER MAY LAST FOR SEVERAL HOURS ... (estimated times of arrival are repeated). 0700 Tsunami reaches Tofino, B.C. 0708 Kodiak reports series of waves: SEA WAVES AT 04352,32 FT AT 05402; 35 FT AT 06302; 9.0 m SEAS DIMINISHING WATER RECEDING. EXPECT 6 MORE WAVES. 0739 First wave arrives at Crescent City. The wave is 0.9 m high. Some evacuees return to danger area. 0750 + Four persons drown at DePoe Bay, Oreg. 0900 Tsunami reaches Hawaiian Islands. Damage slight at Hilo, with three restaurants and a house inundated; at Kahului, a shopping center is flooded.

0920 A 3.7-m wave - probably the fourth - sweeps into Crescent City. This wave, and its successors, destroy or displace more than 300 buildings; 5 bulk gasoline storage tanks explode; 27 blocks are substantially destroyed; there are casualties. 1020 Tsunami reaches east coast of Hokkaido, Japan. 1038 Tsunami reaches northeast coast of Honshu, Japan. 1100 Honolulu Observatory sends final bulletin. It is an all-clear for Hawaii;

29 1 other participants in the SSWWS are advised to assume all-clear status 2 h after their tsunami ETA unless local conditions warrant continuation of alert. 1355 Tsunami reaches Kwajalein. 1910 Tsunami reaches La Punta, Peru. The tsunami does much damage along the coasts of Alaska, Canada, and Crescent City, Calif. Minor damage is reported from the coasts of Washington, Oregon, and the Hawaiian Islands. There are casualties in Alaska, Oregon, and California, but the number is in tens, not hundreds and not thousands. The following is a reproduction from Anon. (1965, p. 24-25) on reminders to population living on coastal areas.

The warning and evacuation of personnel in endangered areas is the job of designated agents participating in the SS W WS. The agent in your location will know what measures to take - and will take them, with your cooperation You can help him, and yourseg by remembering these facts: 1. All earthquakes do not cause tsunamis, but many do; when you hear that an earthquake has occurred in the Pacific Ocean area, stand by for possible communications from your local emergency headquarters. 2. An earthquake in your immediate area should be interpreted as a naturalseismic sea-wave worning; do not stay in low-lying coastal areas afrer a local earthquake has occurred 3. A tsunami is not a single wave, but a series of waves. If you have been evacuated as a result of a SSWWS warning, stay out of the danger area until the entire wave-series has passed 4. Approaching tsunamis are sometimes heralded by a noticeable rising or falling of coastal water. This is nature’s seismic sea-wave warning; it should be heeded by those in low-lying coastal areas. 5. There is at present no way to determine in advance the amplitude, or size, of tsunamis in specific locations. A small tsunami at one beach can be a giant a few miles away; don’t let the modest size of one make you lose respect for alL 6. The SS WWS does not issue false alarms. When a warning is issued, a seismic sea-wave exists. The tsunami of May 1960 killed 61 persons in Hilo, Hawaii, who thought it was “just another false alarm” 7. All tsunamis - like hurricanes - are potentially dangerous, even though they may not strike each Pacific coastline or ab damage at each coastline they strike. 8. Never go down to the beach to watch for a tsunami; when you see the wave you are too close to escape it. 9. Sooner or later, tsunamis visit every coastline in the Pacific. This means that SS W WS warnings apply to you ifyou live in any Pacific coastal area IO. During a tsunami emergency, your local Civil Defense, police and other disaster organizations will try to save your life. Give them your fullest cooperation

Member nations of SSWWS (USA, American Samoa, Canada, Chile, Fiji Islands, French Polynesia, Hong Kong, Japan, New Zealand, Philippines, Taiwan, and Western Samoa) prepared Anon. (1965).

ALASKA Some details of tsunami information dissemination are provided. Tsunami bulletins and warnings are sent to (1) FAA station at Anchorage, (2) U.S. Air Force weather relay facility at Elmendorf Air Force Base, and (3) Alaska Civil Defense, Anchorage. The Civil Defense Headquarters contacts district or area directors of civil defense, mayors, municipal police departments, radio and television stations, newspapers, and wire services. The U.S. weather bureau uses its facilities

292 through Alaskan weather stations located at Barrow, Barter Island, Kotzebue, Nome, Fairbanks, Bethel, St. Paul Island, King Salmon, Cold Bay, Shemya, Anchorage, Cordova, Yakut, Juneau, and Annette, and California, Oregon, and Washington states. Bulletins and warnings are sent via FAA or defense communications system to Hamilton Air Force Base, Calif. This center sends the messages to various places in the three states via the National Warning System (NAWAS). For California, the responsible agencies are the California Disaster Office (CDO) in Sacramento, and the California Highway Patrol in Sacramento. Tsunami alerts are issued usually in the following counties: Alameda, Contra Costa, Del Norte, Humboldt, Los Angeles, Marin, Mendocino, Monterey, Orange, San Diego, San Luis Obispo, San Mateo, San Francisco, Santa Barbara, Santa Clara, Santa Cruz, Solaro, and Sonoma. In Oregon, tsunami bulletins are received by the NAWAS district warning point in Salem or the State Police duty officer. Tsunami warnings are disseminated to counties Clatsop, Coos, Curry, Douglas, Lane, Lincoln, and Tillamook. In Washington, messages are received at the Washington State Patrol radio dispatch office in Olympia. In Hawaii, messages are received at the Honolulu police department dispatch bureau which disseminates the information to the State Civil Defense Agency and Civil Defense agencies in Oahu, Kauai, Maui, and Hawaii. In American Samoa, messages are received via the Defense Communication network by the office of the Governor of Pago Pago. Various US. governmental agencies receive tsunami messages and bulletins. In the Pacific Islands, the U.S. Naval Station at Guam, the US. Naval Station at Midway Island, and the US. Weather Bureau stations at Johnston Island and Marshall Island receive the bulletins. Tsunami messages for Canada are sent to the U.S. Navy Communications in Seattle via the Defense Communications network. From Seattle, the messages are sent to the Canadian Armed Forces naval base in Victoria, B.C. This agency disseminates the warnings to the RCMP and the B.C. Civil Defense. Tsunami bulletins for Chile are sent through the Defense Communications system to Green Belt, Md., and from there, to NASA Minitrack Station in Santiago, Chile. The station relays the information to the Departamento de Navegacion e Hidrografia in Valparaiso, at the same time the messages are sent to the U.S. Navy Communications Station in Balboa (Canal Zone) and from Balboa to Val- paraiso. Alerts are given to various stations and messages could also be sent to Chilean light houses. Tsunami messages for Fiji Islands are sent via FAA Communications to the Aeronautical Communications Station in Nandi and relayed to the Royal New Zealand Air Force Station in Lawthala Bay. From this station, the message is telephoned to the harbor master in Suva. Tsunami messages for French Polynesia are sent via FAA communications to the Aeronautical Communications Station in Nandi, and from there to the Aeronautical Communications Station in Papeete, Tahiti, via radio telegraph. This station notifies the Chief of the Mission hydrographique in Papeete.

293 Tsunami messages for Hong Kong are sent via FAA Communications to the Philippine CAA Communications Station in Manila and then relayed to the Royal Air Force Communications Center in Hong Kong, then the Director of the Royal Observatory is notified. Tsunami bulletins for Japan are sent through the Defense Communications system to the Japan Meteorological Agency. Tsunami messages for New Zealand are sent via FAA network to the Aeronau- tical Communications Station in Nandi, Fiji Islands, or via the US. Navy Communi- cations to Christchurch, N.Z. From Fiji or Christchurch, the messages are transmitted to Wellington and then relayed to the Civil Defense Officer. Tsunami bulletins for the Philippines are sent via the FAA Communications to the Philippine CAA Communication Station in Manila and then relayed to the weather bureau. Tsunami bulletins for Taiwan are sent via the Defense Communications network to Taiwan Defense Command Communications Center in Taipei and then telephoned to the weather bureau. Tsunami messages for Western Samoa are sent via FAA Communications to the Aeronautical Communication Station in Nandi, Fiji Islands, and then relayed by radio telegraph to Faleolo Airport. The information is then sent to the APIA Observatory.

FURTHERDEVELOPMENTS IN PACIFIC TSUNAMI WARNING SYSTEM Murphy and Eppley (1970) discussed developments and plans for the Pacific tsunami warning system since its inception in 1948. Originally it was organized with four seismological stations at College, Sitka, Tucson, and Honolulu, and tide stations at Attu, Adak, Dutch Harbor, Sitka, Palmyra, Midway, Johnston Island, Hilo, and Honolulu. At present all these stations except Palmyra are still part of the warning system. As of October 1969, there were 21 seismic stations and 41 tidal stations participating in the tsunami warning system (TWS). Eleven nations bordering the Pacific are members. The following seismic stations recently joined the TWS: Wellington, Port Moresby, Suva, Tacubaya, Antofagasta, and Easter Island. The inclusion of these will provide quicker and more accurate epicenter determinations in certain areas. The tide stations that plan to join the TWS are Galapagos, Marsden Point (New Zealand), Gambier Island, Tuamotu Arch, Okinawa, Amchitka, Salina Cruz and Manzanillo (Mexico), and Talara (Peru). For locally generated tsunamis in Hawaii, an experimental tsunami warning system is operational. This system consists of seismic and tide-gage stations on the Hawaiian Islands. All data are transmitted to Honolulu Observatory (HO), and recorded there as visual display. Within 10-15 min potential earthquakes capable of generating tsunamis can be identified. The hydraulic gauge data can then be used to supplement the warning. The Alaska Regional Warning System (ARWS) has been in existence in Alaska since 1967. The observatory at Palmer is the center for ARWS, and Adak, one

294 of the Aleutian Islands, is responsible for local warnings within 300 km of Adak TSUNAMIWARNING SYSTEM OF CANADA Dohler (1970) described a tide-gage data telemetry system that has been developed for tsunami warning in Canada. Starting in 1966, Canada installed equipment at three selected stations, Tofino, Victoria, and Langara Island, to supply water-level data, mainly to the Pacific Tsunami Warning System (PTWS) at Hono- lulu, to assist in decision making. Special equipment at the three locations is supplemented with ordinary tide-gage functions. The equipment also has a dialing unit which automatically calls a predetermined party in the event of an unusual water-level condition. The equipment is so instructed that any disturbances in excess of specified conditions, and different from normal tidal oscillations, will be interpreted as a tsunami. Some instrumentation involved in these operations will be briefly described. For details see Dohler (1970). At the Tofino and Victoria stations, the conventional-type tide gages involving a float and a counterweight are used. At the third station on Langara Island a potentionmeter-type, absolute pressure transducer is used. The pressure recording system consists of a pressure head, a three-conductor connecting cable, and a recorder unit. The principle involved here is determination of the height of a water column by measuring the hydrostatic pressure with a resistance potentionmeter-type, absolute pressure transducer. Atmospheric pressure changes that affect the total pressure, measured by the submerged pressure transducer, are compensated in the recorder unit.

TSUNAMIWARNING SERVICE IN JAPAN The Japanese cabinet approved the tsunami warning system on Dec. 2, 1949, and it came into existence on Apr. 1, 1952 (Anon. 1972). The tsunami warning service in Japan consists of three parts: (1) tsunami forecast system - Japan Me- teorological Agency (JMA) is responsible; (2) dissemination system for sending messages to the coastal areas - civil and public organizations; (3) terminal systems to evacuate the people - local authorities. The tsunami warning system of JMA is based on seismic and tidal monitoring stations and analysis centers. If land communication lines are broken for any reason, wireless communication will be used. Figure 6.5 shows the 17 divisions of the Japanese coast, each region under the jurisdiction of one of the five regional tsunami centers. If the epicenter is more than 600 km from the Japanese coast, then JMA Tokyo headquarters has the responsibility of analyzing the earthquake. For nearby earthquakes, the regional centers are responsible. At headquarters two persons are on watch 24 h a day, and at the regional centers one person is on watch round the clock. Japan Meteorological Agency had 106 seismic stations. The four seismic stations at Okinawa joined the JMA network on May 15, 1972, bringing the total to 110. Of these, 69 stations are designated as seismic stations for tsunami warning. Of the 69, 13 are selected for distant earthquakes and are called teleseismic stations for tsunami warning.

295 44 N

Sea of Japan 91

PaciflIC Ocean off

/&inawa

I I I I 1 126 130 138 142 146E FIG.6.5. Japan Meteorological Association network for tsunami warning. Numbers show divisions of coasts assigned to respec- tive tsunami centers; Sapporo, I, 2, 3 ; Sendai, 4, 5; Tokyo, 6, 7, 8, 9, 10; Osaka, II, 12, 13, 14, 15; Fukuoka, 16, 17; e regional tsunami centers; @ national tsunami center. (Japanese National Report)

When an earthquake occurs, a seismic station immediately interprets the seismogram and sends the results by telegram to its tsunami center. This telegram is referred to as “seismic-telegram.’’ Japan Meteorological Agency operates 53 tide-gage stations and 11 are equipped with telemetering devices. When the alarm bell rings at the regional tsunami centers, the location of the epicenter is first determined by the S-P times. The second step is determination of the grade magnitude of the tsunami. The grade magnitude works on a 3-grade scale; grade 1 means major tsunami, grade 2 signifies a minor tsunami, and grade 3 means no tsunami. The grade magnitude is determined by the amplitudes reported in the seismic-telegram (Fig. 6.6). The third step is preparation of the warning message. Five degrees are used for this purpose. In these messages, the tsunami height refers to the amplitude above the normal tide level. First degree means, no tsunami; second degree refers to a tsunami of unknown amplitude; third degree refers to a minor tsunami (not exceeding 2 m); fourth degree refers to a major tsunami (heights from 1 to 3 m); and fifth degree means the warning is cancelled and there is no more danger.

296 200

180

160

140

E' 120 -m U -P g 100 U -m P a d 80

IO 20 30 40 50 EO 70 80s P-S Ih = 40 km) Epicentral Distance FIG.6.6. Japan Meteorological Association tsunami forecasting chart. (Japanese National Report)

For tsunamis originating at distances greater than 600 km from the Japanese coast, the National Center in Tokyo is responsible for warning the public. This warning is based on bulletins issued by the Honolulu Warning Center. Since October 1963, the Soviet Union has been issuing tsunami broadcasts in English from Kha- barovsk. For tsunamis originating in the northern Pacific, Japan uses these broadcasts. Nakayama (1972) described the tide telemetering system employed by JMA for warning of tsunamis and storm surges.

TSUNAMIWARNING SERVICE IN SOUTHPACIFIC The tsunami warning system for New Caledonia and Tahiti is somewhat different from the others and is based on T-phase (see Chapter 3). Actually, this system is applied for Tahiti only because barrier reefs protect the New Caledonia shoreline and no tsunami warning is needed there. This warning system will be briefly described, based on a paper by Talandier ( 1966). Five short-period, vertical seismographs were established in the island group

2 97 of Tahiti and Moorea. The separation between these seismographs varies from 15 to 50 km and they are equipped with filters to eliminate background noise due to swell. Based on the difference in arrival times of seismic waves at these stations, a rough but instantaneous idea of the distance of the earthquake and location is obtained. If two of the five stations report P waves with sufficient amplitude, then an initial alarm is given. Experience shows that any earthquake followed by a tsunami produces large amplitude T waves which travel slower than P waves, but (as sound waves in the water) faster than swell or tsunami. Although the warning time is much shorter than that of Honololu, the Tahiti warning system is much more reliable for that area. However, this statement should not be interpreted as a criticism of the SSWWS, whose purpose is general warning over the whole Pacific area.

TSUNAMIWARNING SERVICE FOR USSR Abouziyarov (1970) described problems of the tsunami warning service in the Soviet Union. The main tsunami threat is from the seismic areas over the continental slope of the deepsea Kuril-Kamchatka trough. Sometimes the tsunamis take only 20-30 min to arrive at the nearest shore. Based on past history, the Kuril-Kamchatka region has been divided into three zones: (1) the most violent tsunami zone on the east coast of Kamchatka between Cape Lopatka and Cape Kamchatka, and the Komandor Islands; (2) moderate tsunami zone opposite the South Kuril Islands beginning at Urup; (3) no tsunamis in the area opposite the middle Kuril Islands (i.e. between the Straits of Krwenstern and the Strait of Boussol). Two types of tsunami sources exist in the far eastern region of USSR. The sources at depths of 5-6 km near the deepsea trough at the positions of the longitudinal breaks could cause large tsunamis affecting distances to 1000 km. On the other hand, sources near the shore associated with lateral faults could cause tsunamis affecting 300-400 km. The tsunami warning service was established following the Kamchatkan tsunami in 1952. Three tsunami warning centers were located at Petropavlovsk-on- Kamchatka, Kurilsk, and Uthno-Sakhalinsk. Table 6.1 summarizes the activities of the tsunami warning service between 1958 and 1965. Special seismographs called UBOPE were developed to quickly determine the epicenters of earthquakes in the magnitude range of 7.0-8.5 and at distances of 150-2000 km. The Kuril-Kamchatka coast is divided into six regions, according to the predicted intensity of tsunami (using statistical relations based on past data): (1) Kamchatka and Kronoki Bay; (2) the Cape of Shipunsky; (3) southern Kamchatka; (4) northern Kuril; (5) Rashuwa, Ketoi, and Simushir Islands; and (6) southern Kuril. SOCIOLOGICALPROBLEMS Few scientific studies have been made on human behavior during a tsunami warning evacuation and the study by Lachman et al. (1961) points to the necessity of unambiguous warning procedures. The authors conducted a survey on human behavior in Hilo, Hawaii, during the 1960 Chilean earthquake tsunami and dis-

298 TABLE6.1 Activity of the tsunami warning service in the Far East from 1958 to 1965. (Abouziyarov 1970)

Processing time Organization reporting, of seismic data data used (mitt) Tsunami characteristic

Nov. 7, 1958 Sakhalinsk Administration of - moderate tsunami south of Kuril Hydrometeorological Service on Islands basis of broadcast by Japan Meteorological Agency

May4, 1959 seismic station Petropavlovsk 5 weak tsunami area Cape Shipunsky and Bay of Avachinsk

Mar. 20, 1960 tsunami station Uthno-Sakhalinsk 10 no tsunami May 24, 1960 Sakhalinsk Administration of destructive tsunami on Pacific Hydrometeorological Service on coast of Kamchatka and basis of information from a man Kuril Islands on port post duty Sevezo-Kurilsk

July 29, 1960 tsunami station Uthno-Sakhalinsk IO no tsunami

Feb. 13. 1961 tsunami station Uthno-Sakhalinsk 7 very weak tsunami south of Kurd Islands

Oct. 13, 1963 seismic tsunami station IO destructive tsunami on Pacific Uthno-Sakhalinsk coast of Urup and Iturup islands

cussed the consequences of an ambiguous tsunami warning system. Some highlights are (Lachman et al. 1961, p. 1405):

“. .. Despite at least 10 hours of warning, the wave killed 61 persons, injured several hundred more, and completely destroyed an estimated 500 dwellings. A study group was organized by the Hawaii Division of the Hawaiian Academy of Science to objec- tively examine the human element in the disaster. The objectives of the research, subsequently undertaken, were to study the subjective interpretations of the warnings and the resulting behaviour. A questionnaire was prepared to be administered to a cross section of the adult population of the affected area.” Characteristics of the sample - Composition of the sample of 327 individuals in terms of sex, age, race, and education is summarized in Table 6.2. The sample included 28 persons who had lost one or more members of their immediate family in the disaster. Also, 50 persons in the sample (15% of those interviewed) had suffered injuries. Of the 50, 47 had not left their homes, the 3 injured among the evacuees either (i) left but later returned to their homes or (ii) sought safety in another stricken area.

Preimpact period - The Siren Signal: The tidal wave warning siren sounded for a 20-min period, more than 4 h prior to the impact of the wave, yet only 40% of the sample evacuated, and presumably this was the percentage of the entire population of the devastated areas who evacuated. Therefore, the question was

299 TABLE6.2. Sample of human behavior during the tsunami of May 1960. (Lachman et al. 1961)

Total sample Nonevacuees Evacuees

No. % No. % No. %

Sex Male 140 42.8 88 44.7 52 40.0 Female 187 57.2 109 55.3 78 60.0 Age 18-27 59 18.0 29 14.7 30 23.1 28-37 76 23.2 41 20.8 35 26.9 3847 77 23.5 45 22.8 32 24.6 48-57 64 19.6 45 22.8 19 14.6 58-67 33 10. I 25 12.7 8 6.2 68 or older 18 5.5 12 6.1 6 4.6 Race Caucasian 8 2.4 3 1.5 5 3.8 Filipino 33 10. I 23 11.7 IO 7.7 Hawaiian 74 22.6 23 11.7 51 39.2 Japanese 178 54.4 123 62.4 55 42.3 Portuguese 15 4.6 14 7.1 1 0.8 Other 19 5.7 11 5.5 8 6.1 Education Grade, intermediate school 151 46.2 100 50.8 51 39.1 High school 152 46.4 81 41.1 71 54.6 College 24 7.3 16 8.1 8 6.2

TABLE6.3. Various interpretations of the siren from individuals interviewed following the tsunami of May 1960. (Lachman et al. 1961)

~__ Total sample Nonevacuees Evacuees

Interpretation No. % No. % No. %

Alert 14 4.8 IO 5.9 4 3.3 Warning 13 4.5 8 4.7 5 4.0 Preliminary signal preceding evacuation signal 71 24.4 55 32.4 16 13.2 Evacuation signal 84 28.9 10 5.9 74 61.2 Signal to await further information 26 8.9 24 14.0 2 1.7 Signal to make preparations 18 6.2 12 7.1 6 5.0 Subjective meaning not ascertainable 65 22.3 51 30.0 14 11.6

Total 291 100.0 170 100.0 121 100.0 posed. “Did you hear the 8:30 siren on Sunday, May 22?” Of those interviewed, 309 (95%) said they had heard the siren; 18 (5%) said they had not. Of those who heard the siren, 127 (41%) evacuated and 182 (59%) did not. The 309 individuals who had heard the siren were asked further if they knew what it meant. Only 18 individuals (6%) said they did not. However, in the course

300 of analyzing the data, it became obvious that not all the 291 persons who said they knew what the siren meant had the same understanding of its significance. Table 6.3 summarizes the various meanings the siren had for the 29 1 individuals who heard it. As the siren meant so many different things to different people, one wondered just what it was meant to signify officially. Consulting the telephone directory (presumably the official medium for disseminating this information), the siren signal was characterized as an “alert,” with no indication as to what behavioral response was expected of the public on hearing it. Responses to the siren are shown in Table 6.4. The disaster victims are classified according to their immediate response to the siren - whether they continued their normal routine, evacuated, or waited for information and instructions. Table 6.5 lists sources of information other than the siren.

TABLE6.4. Relation between ultimate evaluation and immediate response to the siren. (Lachman et al. 1961)

Total sample No n evacue es Evacuees

Immediate response No. % No. % No. %

Did nothing (continued normal routine) 44 15.0 40 23.3 4 3.3 Evacuated 94 32.0 12 7.0a 82 67.2 Waited (for advice, information, etc.) 131 44.5 100 58.1 31 25.4 Other (returned home, etc.) 25 8.5 20 11.6 5 4.1

Total 294 100.0 172 100.0 122 100.0

a Represents individuals who evacuated when they heard the siren but returned home prior to time of impact.

TABLE6.5. Information source concerning the May 1960 tsunami, other than the siren. (Lachman et al. 1961)

Individuals reporting

Information source No. o/c

Relatives, friends 45 17.2 Radio, TV 178 68.2 Government (police, firemen, civil defense) 8 3. I Radio, TV, relatives, friends 22 8.4 Radio, TV, government 6 2.3 Relatives, friends, government 1 0.4 No answer 1 0.4

Total 261 100.0

30 1 6.4 Acoustic and Internal Gravity Waves in the Atmosphere

STRUCTURE OF THE ATMOSPHERE’ Dobson (1963) gave a comprehensive view of the atmosphere in his book; the frontispiece depicts its structure (Fig. 6.7). The nomenclature of different regions of the atmosphere is confusing at present because too many names are used to designate the same region. One important characteristic of the atmosphere is the change of temperature with height. As the temperature reverses several times with height there are three warm and two cold regions. The warm regions are near the earth’s surface, at a height between 40 and 60 km and above 150 km (more or less the top of the atmosphere). The first cold region extends from about 10 to 35 km, whereas the second is around 80-90 km. The exact distribution of temperature with height depends on latitude and to some extent on time of year. Figure 6.7 shows that the temperature decreases upward from the earth’s surface as far as the level called the tropopause. The atmosphere below the tropopause is the troposphere and the region above the tropopause is the stratosphere. Deductions from meteor observations showed that the temperature at about 50 km is approximately equal to that at the surface of the earth. A small amount of ozone is present in the atmosphere, and because the ozone is opaque to ultraviolet light it absorbs solar radiation, and this creates the warm region around the 50-km level. The air generally conducts electricity and conductivity above the 80-km level is much greater than at lower levels, especially during sunlight hours. This region of the atmosphere, called the ionosphere, is basically responsible for allowing radio waves to propagate over great distances. Electrical currents of the region are related to variations of the earth’s magnetic field. At least four regions exist in the ionosphere, known as the D, E, F, , and F, layers, with heights above the earth’s surface varying roughly from 70 km to 400 km. Ionization (when a molecule either gains or loses an electron) in the lower layers of the atmosphere is mainly produced by cosmic radiation. The intensity of ionization in the ionosphere is much greater than in the troposphere or stratosphere, and has an entirely different origin. The higher levels of the atmosphere contain little air and absorb only small amounts of solar short-wave radiation; consequently the ionization is small. When solar radiation reaches lower levels, the abundance of air absorbs the radiation strongly, producing greater ionization. Because of absorption when passing through these layers, the intensity of solar radiation is reduced greatly, and little ionization takes place further down. Thus, there will be some level where the rate of ionization is a maximum. Another important difference between ionization of air in the ionosphere and that in the troposphere and lower stratosphere is that ions are found in abundance in the lower regions, whereas electrons mainly give the ionosphere its characteristics.

ACOUSTICAND INTERNALGRAVITY WAVES IN THE ATMOSPHERE Two major mechanisms appear to generate waves in the atmosphere (Hines 1960): one is tidal oscillations of the atmosphere, and the second is tropospheric

302 10,000 -

Van Allen Radiation Belts 5,000 -

1,000 Ekq'sp here Spray Region Sunlit Aurora

Very Hot Region

Thermosphere Ionosphere 200 - Aurora -

E&rest\ 5 -k

1- Siretus Cloud ./ -100 -50 +50 +loo Temperature (CJ FIG.6.7. Vertical structure of the atmosphere. (Dobson 1963) wind systems. In considering any wave motion in the atmosphere, both gravitational and compressional forces must be considered. Hines showed that even a simple model of the atmosphere (whch is stationary and of uniform temperature and composition) can produce two distinct classes of wave motion with one feature in common; in the absence of dissipative forces both are unattenuated in the horizontal direction. However, in their vertical behavior they differ drastically. The class identified as surface waves, although having an exponential variation vertically,

303 cannot support any phase propagation in the vertical direction. The class identified as internal waves can support significant phase propagation in the vertical direction. The internal atmospheric waves are of two types (Obayashi 1963). In the high-frequency range they are acoustic (sound waves in the atmosphere), whereas in the low-frequency range they are internal gravity waves. Low frequency refers to a period of several minutes when gravitational forces introduce a profound anisotropy. Two characteristic frequencies separate the two limiting cases. One is the Brunt-Vaisala frequency, wg, also called the gravitational stability frequency, and the other is the modified acoustic stability frequency, a,. For an atmosphere in hydrostatic equilibrium, the two frequencies are given by:

and 2 - (L "r=Yg (1 + - 2p az 4~

Here g is acceleration due to gravity, ~(z)is the density of the atmosphere, c(z) is the speed of sound, Y is the ratio of specific heats, and H is the scale height (scale height, H, is the equivalent height of a homogeneous, isothermal atmosphere and is given by -KO* where K is Boltzmann's constant, eA is the absolute temperature, mg g is gravity, and m is the mean molecular mass) of the atmosphere (a function of the vertical coordinate, z). The phase-velocity, V, of the waves is given by:

where

and 2 y (%) and 6 is the angle between the field of gravity and the wave normal. Here o is frequency and T = 27r/o is the period. Waves with periods less than Tu are acoustic waves; those with periods greater than Tg are gravity waves. In the period range, Tu < T < Tg, no internal atmospheric waves exist. For acoustic waves (with periods less than To), the phase velocity is approximately equal to the velocity of sound but tends to 00 as the periods approach Tu. For gravity waves, the phase

304 as well as the group velocities are zero at T = T,, and the velocities increase with increasing T and approach -Y sin 6 for longer periods. The acoustic waves more X or less propagate isotropically whereas the gravity waves propagate mainly in a horizontal direction and exhibit strong anisotropy.

HINES’THEORY FOR ACOUSTICAND GRAVITYWAVES Hines (1960) developed a theory in connection with his interpretation of the irregular motions in the upper atmosphere in terms of internal atmospheric gravity waves. He used a simple model of the atmosphere and brought forth some important features. In the absence of waves, his model at.mosphere is stationary and has uniform temperature and composition. As all three assumptions are drastic, further improvements of the theory are considered. Temperature varies in a complex manner in the atmosphere, as opposed to an isothermal state. The atmosphere, of course, is not stationary; an important force is the Coriolis force due to earth’s rotation. The assumption of uniform composition, although far from the truth, nevertheless is not as drastic as the others for this particular problem. Hines assumed that the superimposed wave motion on the atmosphere has only perturbation magnitude (i.e. second and higher order terms for the amplitude of oscillation in the equations of motion are ignored) and occurs adiabatically. Forces taken into account are gravity, inertia, and pressure gradients. The gravity field is assumed to be constant in magnitude as well as direction. Details of the mathematics leading to the existence of acoustic and internal gravity waves is omitted here but can be found in Hines (1960). Next Hines examined the problem of dissipation. The two basic causes of energy dissipation are molecular viscosity and thermal conduction. Hines also considered the nonlinear effects and concluded that in certain parts of the spectrum and at certain heights, the nonlinearities could be important. FURTHERTHEORETICAL DEVELOPMENTS Although Hines’ original work involved only an isothermal atmosphere, subsequent models used many layers to model realistically the variation of temperature with height. For example, Harkrider (1964) used 40 layers, each a thin isothermal layer. Many numerical studies have been made for the wave-guide modes of acoustic and internal gravity waves for different atmospheric models, all more or less derived from the Air Research and Development Command (ARDC) standard atmosphere (see Fig. 6.8). For periods ranging from a few seconds to a few minutes, these studies were satisfactory in the sense that the calculated values agreed well with observed values. Several authors introduced wind structure and considered its effects on the gravity wave. The work of Weston and vanHulsteyn (1962) is typical. They showed that wind increases the phase velocity of the gravity wave mode. Some important results of the theoretical studies are: in the long-period wave range of 10-100 min, propagation is mainly in the form of gravity waves. There has been some confusion in the atmospheric propagation problems because the distinction between acoustic

305 1

c 8 .-m at I

0 180 240 360 Temperature I'k) FIG.6.8. Standard and extreme ARDC atmospheres. (Harkrider 1964) modes influenced by buoyancy effects, and internal gravity wave modes influenced by compressibility effects, has not been clear in the commonly used term acoustic-gravity wave propagation. Actually, as far as the lower 100 km or so of the atmosphere is concerned, the behavior of gravity and acoustic modes is sufficiently different that no confusion arises. The difference in behavior could be easily visualized in terms of the two limiting frequencies already introduced; the acoustic low-frequency cutoff, ma, and the Brunt-Vaisala frequency or the internal gravity wave, high-frequency cutoff, wg.In the literature several notations have been used for these; Tolstoy and Pan (1970) used woand N. In the regions where w, > w ,there is no ambiguity in the two types of modes. However, centered at about 120 fm in the atmosphere there is an anomalous zone in which wg > w,.This anomalous zone is probably important for ground-level detection of surface gravity waves excited by atmospheric nuclear explosions as will be shown later. Tolstoy and Pan (1970) showed that simple models involving two to four layers in the atmosphere, can be used to determine the propagation and dispersion of the lowest two to three gravity modes of the atmospheric wave guide. Two important facts brought out by this paper are: these simple models could explain the long- period portion of the 300 m/s group velocity plateaus in terms of coupling effects of the internal gravity wave guide. Tolstoy and Pan cautioned that, as this result is is obtained from an atmospheric model with two incompressible layers, the 300 m/s

306 value is somewhat fortuitously related to the velocity of sound in air. Also, surface gravity-wave modes of the atmosphere are better understood by this approach. Another concept introduced pertains to the top boundary condition in the atmosphere. For example, Pfeffer and Zarichny (1962), Press and Harkrider (1962), Harkrider (1964), Daniels (1967), and Harkrider and Wells (1968) used a free-surface condition, at several heights arbitrarily, to compute surface modes of the atmosphere, but did not explain the logic of this. Tolstoy (1967) showed that it is indeed impossible to justify the use of a free-surface boundary condition, and thereby make the selection of H, the effective atmospheric wave height, less arbitrary. For this he used the concept of vacuum, i.e., at heights where the mean free path,Q, of the neutral gas molecules exceeds the wavelength, L, of the disturbance, the medium is considered to be vacuum. Hence, for heights where Q << L, the usual continuum equations apply. Then the effective free surface of the atmosphere is at height, z = H, where R = L, This means the effective height of the atmosphere depends on the wavelength of the disturbance. However, the region, Z = H, at which Q 2 L is not infinitesimal, but is a transitional layer of thickness, d (say). In this transition layer, strictly speaking, Boltzmann’s equations must be applied. For d << L, the layer mainly behaves as a region with large viscosity and strongly attenuates waves with periods less than 10 min. Another consideration is that (at the heights relevant here) because of large ionization, hydromagnetic interactions could occur. However, Dungey (1954), Fejer (1960), and Hines (1955) showed that these interactions are effective only for periods greater than or equal to 3 h. Hence, it can be assumed that the atmosphere acts as a window for surface gravity waves, with periods between 10 and 200 min.

TOLSTOYAND PAN’SMATHEMATICAL MODEL The simple mathematical model of Tolstoy and Pan (1970) for wavelengths L > 300 km will now be discussed, with particular attention to aspects of the propagation directly relevant to recent observations with microbarographs. They showed motions with 600 m/s and energy concentrated around the 15-min period. The logic behind using a simple model is given by Tolstoy and Pan (1970, p. 35): “Since in the present study we are interested chiefly in internal and surface gravity modes with periods > 10 min and wavelengths > 200 km, a small number of layers provides adequate approximations for discussing the propagation properties of atmospheric waves. One must remember that the frequency-wave number relationship and the group velocity may be written as quotients of two quadratic forms in the wave amplitudes (Boil, 1957). These are stationary with respect to variations of the amplitude, so that in most calculations relatively large errors in the amplitude distribution of the displacement field can be tolerated, without appreciably affecting the eigenvalue and propagation velocity calculations. Thus, even though the amplitude of the vertical displacement predicted by an oversimplified model may be in error, the characteristic curve calculations can be quite accurate. “We have limited ourselves to calculations on two and four-layer models, illustrat- ing the effects of compressibility, of the upper boundary condition, of the earth’s

307 rotation, and of layering upon the propagation of long periods. Within the limitations imposed by the small number of layers, we have used models that represent fits to the Vaisala frequency, the density, and sound velocity functions in the earth’s atmosphere.” Tolstoy and Pan’s model allows interpretation of recent observations by microbarographs. Tolstoy and Herron (1970) also gave a new interpretation of the traveling disturbances in the ionosphere due to nuclear explosions. This is possible because in the moderately long-period range of 10-30 min there is considerable separation between the velocities of surface and internal gravity waves. The main difficulty appears to be in the magnitude of the pressure signal. Harkrider and Wells (1968) showed that the observed pressure amplitudes from peak to trough of 10-100 pb would require unrealistically large vertical displacements of the atmosphere above 100 km. Other calculations with simple models essentially confirm their result. Tolstoy and Pan (1970) attributed this to ignoring the anomalous zone between 110 and 200 km, and ignoring the influence of the winds at high altitudes. When these are taken into account the difficulty disappears. That is, the vertical displacements of the atmosphere above 100 km need not be ridiculously large to be consistent with ground-level pressure amplitudes in the range of 10-100 pb, Next Tolstoy and Pan considered the question of the attenuation of the surface gravity waves and showed that generally the attenuation is small for periods greater than 10 min for the modes m = 0 and 1. Liu and Yeh (1971) studied the excitation of acoustic-gravity waves in an isothermal nonrotating atmosphere; because their model is highly idealized it is not included here. Yeh and Liu (1972) gave a good review of the propagation of waves in the ionosphere.

6.5 Atmospheric Nuclear Explosions Before discussing the acoustic-gravity waves that could be generated by nuclear explosions in the atmosphere, some general effects of atmospheric nuclear explosions are described. The release of a large amount of energy in a small volume in a short time interval is referred to as an explosion (Glasstone 1962). This quick release of energy, whether from conventional chemical explosives or from nuclear materials, produces a great increase in temperature and pressure, so that all materials in the immediate vicinity are transformed into hot and compressed gaseous forms. At high temperatures and pressures these gases expand rapidly and cause a pressure wave, generally called a shock wave. However, this term is used only for underground and underwater tests, whereas for explosions in the atmosphere the term used is blast wave. Two factors primarily determine the amount of energy (in proportion to the total energy) that will arrive as thermal energy at a given point at some distance from the site of the explosion; the nature of the nuclear device and the nature of the surroundings. At higher levels in the atmosphere there is little air to interact with the energy of the explosion and the proportion of the fission energy converted into the blast wave is small in comparison with the portion that goes into thermal

308 energy. On the other hand, at a relatively low level of 30.5 km, about 50% of the energy goes into blast, 35% into thermal energy, 5% into the initial nuclear radiation, and the rest (Le. 10%) is accounted for by residual nuclear radiation. In the extreme case of an underground explosion that is confined, little thermal radiation escapes into the atmosphere. Glasstone (1962) identified five types of explosions: air burst, high altitude burst, underwater burst, underground burst, and surface burst. He used the differ- ences in the nature of the accompanying fireball to set them apart. The fireball is a hot and luminous mass (roughly spherical) created from the residues of the exploded nuclear device as well as from the surroundings. An explosion at a level in the atmosphere not greater than 30.5 km, followed by a fireball that does not touch the ground surface at the time of its maximum brilliance, is called an air burst. Glasstone stated that a 1.02 X lo9 kgm device would cause a fireball with a radius of 0.88 km. If the device is exploded at an altitude greater than 30.5 km it is called a high-altitude burst. A surface burst is when the weapon is exploded either at or just above land or water. If the explosion occurs beneath the ground, it is called an underground burst, and beneath water it is called an underwater burst; these can be called subsurface bursts. Certain terms will now be defined following Glasstone (1962). The overpressure in the blast wave is the excess over the atmospheric pressure of 101.4 kN/m2 (6.895 kN/m2 = 1 psi) at standard sea-level conditions. Most material damage from a nuclear explosion is caused by the shock or blast wave. A difference of air pressure on different surfaces of a building, for example, will produce a force on the building. The so-called overpressure should be considered when estimating the destructive effects of a blast wave. The overpressure is maximum at the shock front and this value is called the peak overpressure. As the blast wave moves away from the origin, the overpressure at the front decreases steadily and after a short time (when the shock front has traveled some distance from the fireball) the pressure behind the front becomes less than that of the atmosphere. This is the so-called negative phase of the blast wave. That is, in this region the air pressure is less than the ambient atmospheric pressure, and an underpressure exists. Another important parameter is the dynamic pressure, proportional to the square of the wind velocity and the density of the air behind the shock front. EFFECTSOF NUCLEAREXPLOSIONS ON ATMOSPHEREIONIZATION Ionization is the phenomenon when ion pairs of separated electrons and positive ions are formed in the ionosphere. There are mainly three regions in the ionosphere - D, E, and F regions. In the daytime, and especially in summer, the F layer splits in two, Fl and F2. The electron density is a maximum in each region. However, the electron density increases generally with altitude, i.e. it is greater for the F region than for the E region, which has greater electron density than the D region. In the lower part of the atmosphere, i.e. below 48.3-56.3 km, the air is dense and the probability of collisions between free electrons and atmospheric atoms and molecules is great. Because of this, electrons are rapidly attached to the neutral particles and ionization cannot be produced. Even if it is produced, it is destroyed

309 immediately because, on the average, the life span of a free electron in this part of the atmosphere is less than a microsecond. In the ionosphere, electrons and ions are produced by interaction of solar radiation with atoms and molecules of the atmospheric constituents. They are destroyed by combining either with neutral particles or positive ions. Of the two destructive processes, the former dominates at lower levels where air density is greater, whereas the latter dominates at higher levels where electron density is greater. The effect of a nuclear explosion on atmospheric ionization is mainly because of an increase in electron density in the surrounding region. The added electrons can affect all electromagnetic communication, either by attenuating the signal or by changing its direction of propagation through refraction.

ACOUSTIC-GRAVITYWAVES FROM ATMOSPHERICNUCLEAR EXPLOSIONS Donn and Shaw (1967) studied the atmospheric nuclear tests conducted by the USA and USSR during 1952-62 and, based on data recorded by a global network of stations maintained by the Lamont Geological Observatory of Columbia University, they arrived at interesting conclusions about pressure waves due to these explosions. In fact, effects of the high-yield atmospheric nuclear explosions are comparable to those of the enormous Krakatoa eruption of 1883 (see Chapter 2) and the impact of the Siberian meteor in 1908, when pressure waves in the atmosphere traveled round the globe more than once, with appreciable amplitudes. Donn and Shaw used the data from sensitive microbarographs at 15 recording stations, and acquired 208 records of 45 nuclear explosions. The dispersion relations of the acoustic-gravity waves in the atmosphere were used to analyze the records and the authors concluded that: “The initial spherical wave at the source, which is modified into a cylindrical wave by the layered structure of the atmosphere, is composed of a broad spectrum of pressure waves whose frequencies, which propagate away at about the speed of sound in air, range from audible sound to about 0.002 cps. At distances of a thousand or more kilometers, the spectrum becomes considerably narrowed, with the highest frequency detectable being about 0.03 cps (30 sec in period). It is more convenient to refer to period rather than frequency for these infrasonic waves. Such waves are also referred to as acoustic-gravity waves, since their propagation characteristics are controlled by both gravity and the acoustic properties of the atmosphere. “Because the atmosphere is a dispersive medium for the long waves within this period range (30 to 500 sec), the initial impulse becomes dispersed into a train of waves. The exact character of the period-velocity dispersion is a function of the temperature and wind stratification of the atmosphere along the propagation path.” Daniels et al. (1960) studied vertically traveling shock waves in the ionosphere caused by nuclear explosions at the surface of the earth. They interpreted the distortions in ionosphere recordings as the result of retarded sound waves. Daniels and Harris (1961) reinterpreted the record (because the velocity of tKe shock wave was 115 m/s) as an ordinary hydrodynamic shock wave. They also mentioned that the pressure amplitude of this upward traveling shock wave was large enough to be detected by acoustic methods at ground level.

3 10 Webb and Daniels (1964) reported ionospheric oscillations following the Soviet atmospheric nuclear test of Nov. 1, 1962. They measured the rotation of the polarization plane of the radio waves during their travel through the ionosphere, and this parameter was proportional to the electron density. They used a 151-MC/S transmitter at Fort Monmouth, N.J., to reflect signals from the moon. The signals were received at the University of Illinois at Danville, between 1100 and 1800 CST on Nov. 1, 1962. Thus, the signals traveled twice through the ionosphere. From newspaper and other reports it was known that the Soviet Union had conducted a nuclear test in the atmosphere a few hours earlier. The record at Danville showed an oscillation with 30-min period, larger than the period recorded by ground-level microbarographs. Webb and Daniels attributed the oscillation to acoustic waves in the atmosphere caused by the explosion. These acoustic waves are expected by theory to have a group velocity of about 730 m/s, but the authors were unable to check this from the oscillations that may have started before the record began. Long-period ionospheric oscillations have been known to continue for several hours following a nuclear explosion (Daniels and Harris 1958). Tolstoy and Herron (1970) studied the atmospheric gravity waves excited by nuclear explosions during 1967-68. The data used were from the large aperture, i.e. a 250 km x 200 km array of 6-12 low-frequency microbarographs with period pass-band 1-60 min, in the New York-New Jersey area. The spectra of these gravity waves showed peaks at 15-min periods and the average group velocity was approximately 600 m/s. The authors deduced that these were indeed atmospheric surface-gravity waves, based on agreement of the dispersion and attenuation of the recorded waves with expected theoretical relations. Tolstoy and Herron (1970, p. 59) concluded that the 600m/s, 15-min period arrivals were probably surface gravity waves traveling along what is effectively the top of the atmosphere. The height, H, of this “effective free surface” depends on the wavelength, L, as it corresponds approximately to the region where the mean free path, Q, of air molecules is of the order of L (indeed, the condition explicitly applied in making calculations is 2nP/L = 1). They thought that a number of published ionospheric observations of fast- traveling disturbances generated by U.S. and USSR thermonuclear tests in the early 1960s could probably be explained in a similar way. It appears that reported ionospheric disturbances with horizontal group velocities 2 500 m/s are surface gravity waves, whereas lower velocities correspond to internal gravity waves (Hines 1967). Acoustic modes of propagation (Wickersham 1966) are possible for spectral components with periods shorter than 10 min. Earlier, the problem of the pressure amplitude of atmospheric gravity waves was mentioned. To be consistent with the pressure amplitudes at the ground, it appeared at the outset that unrealistically large vertical displacements of the atmosphere above 100 km altitude would be needed. Later the problem was explained as neglect of an anomalous zone. Tolstoy and Herron (1970, p. 60) proposed that: “A much oversimplified one-layer model of the atmosphere (Tolstoy, 1967) can be used to give peak-to-trough pressure perturbations at ground level of 2 pb km-’ displacement at these heights, a result which would be quite adequate were it not

311 for the fact that a more realistic two-layer model decreases this estimate to 0.2 pb or less. Proper inclusion of the effects of a layer between altitudes of 110 and 150 km, in which the acoustic cut-off frequency of the medium is lower than the Vaisala frequency, might eliminate the discrepancy; the problem is discussed in more detail elsewhere (Tolstoy and Pan, 1970).” Hines (1967) clarified some confusion that arose in connection with the ionospheric disturbances caused by the Soviet upper atmospheric nuclear test over Novaya Zemlaya on Oct. 30, 1962. Obayashi (1962, 1963) attributed these perturbations in the F-layer critical frequency to atmospheric surface gravity waves with periods in excess of 10 min. Wickersham (1966) differed and proposed that the ionospheric disturbances were due to fully ducted acoustic-gravity waves in the atmosphere. Hines showed that Wickersham’s interpretation is not well founded and Obayashi’s original interpretation is correct.

6.6 Atmospheric Disturbances Generated by Earthquakes and Volcanic Explosions Benioff et al. (195 1) presented data on atmospheric sound waves generated by an earthquake. A microbarograph at Pasadena recorded a train of waves with periods ranging from 5 to 1 s, attributed by the authors to an earthquake that occurred Jan. 24, 195 1. Donn and Posmentier (1964) studied the ground-coupled air waves from the Alaskan earthquake of March 1964, by the micropressure fluctuations recorded on microbarovariographs at Honolulu, Berkeley, and Palisades. The microbarovariographs have a sensitivity in the range of a few to several hundred microbars and a flat response up to 300 s. The authors indicated that atmospheric pressure oscillations due to earthquakes may be produced by three mechanisms in the region directly surrounding the recording station, the epicentral region, and a region where there is resonant ground-air coupling. Rayleigh waves emanating from the epicentral area can generate large pressure fluctuations in the atmosphere by impulsive effect, provided their vertical displacement is large. The following relation holds between the air- pressure perturbation, p; density of air, p; velocity of sound in air, c; and velocity of the vertical ground motion, (Donn and Posmentier 1964):

p = pcv (6.10)

The formula assumes that the interference from sound waves in adjacent regions is negligible. This condition is met, provided resonant coupling between ground and air is not appreciable and wavelengths of the Rayleigh waves are much greater than the height of the receiver (microbarograph) above the ground. If p = 7.19 X l0y3 g/cm and C = 330 m/s, the above equation becomes: a p = 247 7 (6.1 1) wherep is peak-to-peak pressure in microbars, a is the double amplitude of ground motion in centimeters, and t is the period of motion in seconds.

312 In the second region the vertical displacement of the ground near the epicenter will generate atmospheric acoustic waves. Donn and Posmentier ( 1964) visualized this mechanism as somewhat analogous to generation of acoustic-gravity waves from nuclear explosions. In the third region, vertical displacement of the earth’s surface far from the epicenter might generate acoustic waves that could be detected at distances of a few hundred km, provided the appropriate conditions for ground-air coupling exist. It is this type of mechanism that accounted for the ground-coupled air waves reported by Benioff et al. (195 1) for the Imperial Valley earthquake of Jan. 24, 195 1. Donn and Posmentier estimated from the seismograms that the maximum vertical ground motion at Palisades was 4.2 cm, attained by the initial Rayleigh wave. The observed maximum pressure fluctuation at Palisades was 40 pb, and from Equation (6.1 1) (assuming a period for local vertical ground motion as 23 s, the period of strongest pressure waves) one gets 3.72 cm for the required vertical motion. Because this value approximates the observed value of 4.2 cm, it was concluded that the main source for the pressure fluctuations was indeed the local vertical ground displacement due to the earthquake. Donn and Posrnentier concluded that dispersion curves for Palisades and Honolulu showed typical continental and oceanic Rayleigh waves, respectively, and the record for Berkeley showed both; mechanism 2 seemed to have worked only at Berkeley. The Berkeley record showed large-amplitude, long-period waves in the period range of 3-5 min. The superimposed smaller waves were continuations for earlier ground-coupled air waves. The structure of the initial long-period waves is different from that of the background gravity waves. Donn and Posmentier attribute the dissimilarity between these waves and those caused by nuclear explosions (those due to earthquakes show clearer dispersion) to the fact that in earthquakes the generation area is large, whereas an explosion is a point source.

10NOSPHERIC EFFECTSDUE TO EARTHQUAKES Davies and Baker (1 965) detected ionospheric disturbances associated with the Alaskan earthquake of March 1964. They made observations at Boulder, Colo., on frequencies of 4 and 5 Mc/s with vertical propagation, and on 10 Mc/s at WWVH in Hawaii, 5000 km from Boulder. During 4 years of observation, they recorded only one other such incident; that was due to the July 9, 1962, nuclear explosion at Johnston Island.

SOVIET WORK ON VERTICAL PROPAGATION OF ACOUSTICWAVES DUETO AN EARTH- QUAKE Important contributions were made by Romanova (1.970) and Petukhov and Romanova (197 1) to understand generation and vertical propagation of acoustic waves in the atmosphere due to an earthquake, and subsequent heating in the atmosphere at different levels, caused by these waves. The earlier paper deals mainly with development of the theory, the later paper with its application. Results of the second paper are summarized here. Petukhov and Romanova applied Romanova’s theory to calculate the dissipation of infrasonics from earthquakes in the atmosphere in the form of heat. In

3 13 particular they estimated the rate and amount of atmospheric heating due to acoustic waves traveling upward from two great earthquakes, the Alaskan earthquake of 1964 and the Katinoko, Japan, earthquake of 1968. This atmospheric heating had a significant role in the observed ionospheric disturbances. The authors essentially started with the same initial conditions as Donn and Posmentier (1964), i.e. with Equation (6.10) or (6.1 1). This means the Rayleigh waves from the epicenter propagate over the earth's surface with a velocity of roughly 3 km/s, which is supersonic for air. The vertical motion of the Rayleigh waves produces a vertical impulse effect in the atmosphere, and produces pressure perturbations. The relation between the pressure perturbation, p, and the displacement, a, of the ground surface is given (as in Donn and Posmentier's work) by (6.10). Under the linear theory for small amplitude oscillations, the Stokes-Kirchoff equation for the attentuation coefficient, (Y, of acoustic waves is given by:

(6.12) where w is the frequency expressed in Hertz, p is air density, Y is ratio of specific heats, pa is the molecular weight of air, n is density of air molecules, and P, is the Prandtl Number (0.7 for air). The equation used for atmospheric heating is:

(6.13) where 0, is the absolute temperature, E is the wave energy, c is the velocity of sound, and bar denotes averaging over a period. After some algebra (6.13) becomes:

(6.14) where cp is the specific heat at constant pressure. This formula was used to calculate the temperature change for the atmosphere over Berkeley, Boulder, and Palisades for the 1964 Alaskan earthquake and for the atmosphere over Maui for the 1968 Katinoko earthquake. The calculations showed that maximum heating occurred in altitudes of 170-190 km. At this altitude, the atmosphere over Berkeley was heated by 900". Although this value appears large, the peak-to-peak amplitude of the ground motion with a period of 23 s was 12.6 cm. Ionospheric disturbances were observed associated with this earthquake (Le. 1964 Alaskan earthquake). An important method of identifying these disturbances was an increase of about 20 km in the altitude of the critical reflecting layers. The role of atmospheric heating in this could be significant, because, when the 170- 190 km altitude layer (with maximum heating) expands with heat, it disturbs the layers above it. Under the assumption of an isobaric process, the rise in the

3 14 critical layer z = zo caused by heating and expansion of the underlying layers can be computed from the following relation:

(6.1 5)

Here To(z) is temperature at height, z, in the atmosphere before heating starts, and A& = (ae,/az) At; aGu/az is calculated from Equation (6.14). This calculation gives a rise of 20 km for the layer at 240 km altitude over Boulder and the observed rise is of the same order. Over Maui, the 200-km altitude layer theoretically should rise 15 km, although it was observed to rise only 5-6 km. In contrast to the linear theory, the nonlinear theory shows that lower atmospheric layers are heated more and higher layers are heated less, thus the nonlinear theory gives a smaller rise of the critical layers. Another important point Petukhov and Romanova (1971) emphasized is that because of the decrease of atmospheric density with altitude, temperature fluctuations will mainly travel upward, and a new isothermal condition is rapidly established in the atmosphere above the heat source at the ground. To obtain this isothermal condition at the 170-200-km level would only take about 3 h.

SEISMICWAVE COUPLING TO THE IONOSPHERE Details on analysis of ionospheric disturbances due to seismic waves are presented. Yuen et al. (1969) studied the May 16, 1968, Hachinohe, Japan, earthquake and analyzed and compared the seismic, atmospheric, and ionospheric data. Some important results are: acoustic waves generated in the atmosphere by seismic waves due to the earthquake traveled to 300 km altitude and created oscillatory disturbances in the ionosphere. Acoustic waves in the period range of 20-25 s were attenuated more strongly than those with longer periods of the order of 2min. Value of the results lies in the fact that the observed pressure changes, determined from Doppler records, agreed well with theoretically expected values. Donn and Posmentier (1964), for the Alaskan earthquake of March 1964, could not directly compare the seismograms with microbarograms because the vertical component seismograms were illegible, due to large ground displacements. For the Hachinohe earthquake, the average double amplitude of the Rayleigh wave was 0.28cm, as shown by the seismograph at the Hawaiian Institute of Geophysics (HIG). The average Rayleigh wave period was taken as 25 s, and this gives from Equation (6.1 1) a value of 2.8 pb for the average peak-to-peak pressure variation, an order of magnitude less than the pressure variations associated with the 1964 Alaskan earthquake. The pressure fluctuations for the Hachinohe earthquake were too small for the microbarographs at Honolulu to record. However, Yuen et al. (1969, p. 2258) were able to deduce the pressure changes from the ionospheric Doppler records. They described the HF Doppler technique, where the instantaneously received frequency of a highly stable ionospherically reflected signal is measured. The Doppler frequency shifts indicate changes in the ionosphere caused by movements of the reflecting layer or by variations in the electron concentration below the point of reflection.

3 15 A few minutes after the Rayleigh waves reached Hawaii, short-period oscillations were produced at heights near 200km and long-period oscillations were produced near 300 km. The Faraday rotation of the polarization of CW signals, transmitted from geostationary satellites, were measured to continuously obtain total electron content data. The period after the earthquake showed no unusual behavior. This indicated that no significant net production or loss of electrons were associated with disturbances on the Doppler and ionogram records, and implied that these disturbances were produced by electron concentration pressure waves in the ionosphere. According to Yuen et al. (1969): “An acoustic longitudinal pressure wave with a vertical velocity component would be expected to move the contours of constant electron concentration up and down with an oscillatory motion similar to the pressure wave itself. This would mean that the height of reflection for HF radio signals such as those used ifi the Doppler system would also move up and down with an oscillatory motion. The pressure wave at ground level generated by a Rayleigh wave would be proportional to the seismic motions at the earth’s surface. Thus similar shapes would be expected on the seismic and Doppler records.” To support their contention that the ionospheric disturbances shown on the Doppler recordings were caused by acoustic waves generated by Rayleigh waves from the earthquake, Yuen et al. used ray tracing, and calculated the attenuation by the 1966 U.S. standard atmosphere. In this calculation, the exospheric temperature was taken to be 1200” k over Hawaii at the time of the earthquake. Assuming adiabatic conditions, Yuen et al. (1969) calculated the pressure change from the change in the electron concentration.

ATMOSPHERICDISTURBANCES DUE TO VOLCANIC EXPLOSIONS The effects of the Krakatoan eruption of August 1883, and the atmospheric disturbances on the water-level changes that occurred globally were treated in Chapter 2, but information on the observed atmospheric pressure disturbances is presented here. Scott ( 1883) described the atmosphere pressure fluctuations observed following the Krakatoan eruption, and Strachey (1883) made deductions based on this work. Scott mentioned that the volcanic explosion probably occurred between 1600 on Aug. 26, and 0500 Aug. 27, 1883, local time (in Greenwich time, between 0900 and 2200 Aug. 27). The pressure disturbances discussed by Scott are summarized in Table 6.6. On the average, the pressure wave took 36 h 37 min to travel around the earth from east to west, and 35 h 17 min from west to east. By working back with this data, Strachey deduced that the Krakatoan eruption occurred at 0224 GCT (or 0924 local time) Aug. 27, 1883. The atmospheric pressure wave on the average traveled 674 mph from east to west and 706 mph from west to east.

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